Microclimates

5/19/2018 Microclimates

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MICROCLIMATE

M W Rotach, Swiss Federal Institute of Technology,

Zurich, Switzerland

P Calanca, Swiss Federal Research Station for

Agroecology and Agriculture, Zurich, Switzerland

Copyright 2003 Elsevier Science Ltd. All Rights Reserved.

Introduction

On the global scale the climate is mainly determinedby the incoming solar radiation. The word climateitself stems from the Greek klima, which meansinclination andis appropriate for describing theeffectsof latitude on the availability and distribution of solar

energy. On very small scales, however, global factorsset only the boundary conditions, while the actualclimate at any particular location is defined by thelocal surface properties. An outstanding example isthat of an oasis surrounded by a desert. Obviously, it isnotthe Earths orbital parameters northe compositionof the atmosphere that determine the climatic statewithin the oasis but rather the local soil moistureconditions.

The microclimate of a particular location canhence be defined as the statistical state of theatmosphere in the layer being affected directly bythe characteristics of the underlying surface. In

this sense, microclimatology is the study of thelong-term average and typical variability of climatevariables in this lowest layer of the atmosphere,and physical microclimatology can be thought of asthe study of the processes by which the lowest layerof the atmosphere responds to surface boundaryconditions.

The processes controlling the microclimate at aparticular location involve time scales from a fewseconds to several years. For example, evapotrans-piration from a vegetated surface is regulated on theshort term by the physiological processes of theplants, whereas decadal changes in the species

composition of that surface determine its long-termtrends.Like the temporal scales, the spatial scales usually

associated with the term microclimate encompassseveral orders of magnitude. The microclimate ofanimals or single plants deals with scales rangingfrom centimeters to meters. On the other hand, thecharacteristic scales associated with the microclimateof an extended downtown area with high-rise build-ings are on the order of hundreds of meters, orkilometers.

Variables Describing and Determining

Microclimate

As usual in climatology, the microclimate of a partic-ular location is characterized through a number ofclimate variables. Traditionally, these variables arethose describing the thermodynamic and dynamicstate of the atmosphere, i.e. radiation, temperature,humidity, wind speed, and pressure (density). Depend-ing on the research focus, other variables, such as tracegas concentrations, should be added to the list. Forexample, the health and the comfort of the everincreasing number of people living in cities is directlyrelated to the concentration and distribution of airpollutants, which are additionally required to charac-

terize the microclimatic state of this particularenvironment.

As for the surface properties determining themicroclimate, the following groups of variables arerelevant in the case of simple, flat surfaces: radiativeproperties (albedo, emissivity); aerodynamic charac-teristics (roughness length, zero plane displacement);thermal properties (heat capacity and conductivity),and properties affecting the moisture status (hydrauliccharacteristics of the soil, type of surface cover). Atsites with a more complex geometry, slope andexposition have an important impact, in particularon the components of the radiation balance. These

surface propertiesare notstatic, but reflect forinstanceforcingby thelarge-scale wind field, diurnal changes inthe soil moisture conditions, seasonal changes in thevegetation, or the presence or not of snow.

The temporal and spatial distributions of the statevariables are governed by the conservation equationsfor energy, momentum, and mass (dry air, water vapor,tract gases).Owingto thegenerally turbulentnatureofthe flow in the vicinity of the Earths surface, the termsrepresenting turbulent transport in these equations areparticularly important and receive much attention inresearch of microclimate.

The Lowest Layers of the Atmosphere

The atmospheric layer affected by the local surfaceproperties is called the planetary boundary layer. It hasa characteristic depth of the order of 1000 m and canbe dividedinto the upper or outer layer (the uppermost90%) and the inner or surface layer (that is, the lowest10%), respectively. However, a direct influence of thesurface characteristics on the atmospheric state isobserved only in the lowest part of the surface layer, in

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the immediate vicinity of the roughness elements. If

this layer of influence has any discernible thicknessthen it is because of a nonnegligible vertical extensionof the so-called roughness elements (stones, vegeta-tion, trees, buildings). This layer is therefore usuallycalled theroughness sublayer, and theupper part of thesurface layer is then referred to as the inertial sublayer.Note that over a relatively smooth surface such asshort grass or sand the roughness sublayer becomesvery thin, and the inertial sublayer is often associatedwith the entire surface layer (Figure 1).

On the basis of these considerations, microclima-tology can be thought of as the study of the climaticstate of the roughness sublayer or specific entities

therein, whereas micrometeorology is more properlyrelated to the study of the dynamics and thermody-namics of the surface layeror even the entire planetaryboundary layer.

The Roughness Sublayer

In the roughness sublayer, the flow is affected by theindividual roughness elements, and is hence fullythree-dimensional in nature. The upper boundary ofthe roughness sublayer, zn, is the level at which thehorizontal variability vanishes and the flow becomeshorizontally homogeneous. Properly scaled profiles ofeither mean flow characteristics or turbulence statis-

tics will then merge to one curve that is characteristicfor the underlying surface. The depth of the roughnesssublayer depends on the height and distribution of theroughness elements. For most surfaces 2hozno5hcovers the range of estimates, where h is the averageheight of roughness elements.

The Canopy Layer

The lowest part of the roughness sublayer is the so-called canopy layer. It ranges from thesurface up to the

average height of roughness elements. In this layer,form (or pressure) drag and viscous drag on anyindividual roughness element are both significant andlead to a retardation of the flow. In addition, thematerial and orientation of the obstacles give rise to

large variation in the energy balance and primarily theradiation balance. Moreover, within the canopy layerthere is a substantial variability in the source/sinkdistribution of sensible heat, water vapor, or tracegases.

Figure 2 shows the height ranges covered by theroughness sublayer and its lower part, the canopylayer, in a dimensionless form. Arrows indicate thetypical situation for a surface covered with crop, trees,and houses, respectively. We observe that the verticalrange, to which the study of microclimate is confined,can extend up to several tens of meters or more in thecase of an urban surface (if we use typical values for h).

In the case of a shallow boundary layer (smallboundary layer height zi and large roughness ele-ments), no inertial sublayer may be present at all.

So far we have dealt only with rough but horizon-tally homogenous surfaces like an extended cornfield,a large forest, or a uniform city. In the case of small-scale variability of surface properties, a so-calledinternal boundary layer forms downwind of eachmajor change in the surface characteristics. This

z= zi:

z=0.1zi:

1000m

100m

Boundarylayer

Surfacelayer

Inertial sublayer

Roughness sublayer

Canopy layer

Free troposphere

z 3h

z= h

Height(log scale)

Figure 1 Conceptual sketch and terminology for the lowest

layers of theatmosphere overa roughsurface.Note the logarithmic

height scale. The levelzirefers to the boundary layer height andh

denotes the (average) height of roughness elements.

100

101 102 103 104

101

102

103

104

City

ForestCropz/zi

zi/h

Roughnesssublayer

Canopy

z=h

BL height /Roughness element height

Height/BLheight

Inertial sublayer

z*=3h

Figure 2 Sketch of the vertical extension of the various layers

over rough surfaces and their variation with the nondimensional

quantitiesz=ziandzi=h. The heightzirefers to the boundary layerheight and hdenotes the (average) height of roughness elements.

A value of zn=h 3 is assumed for the height of the roughnesssublayer. The arrows city, forest, and crop are based on typical

values for the height of the roughnesselementshandthe boundary

layer heightzi. (Reprinted from Rotach MW (1999) On the urban

roughness sublayer. Atmospheric Environment 33: 40014008,

with permission from Elsevier Science.)

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internal boundary layer may be thermal (thermalinternal boundary layer) if there is primarily a changein the thermalproperties of the surface, as for exampleon a shoreline. Alternatively, if mainly the roughnesschanges then the internal boundary layer is called

mechanical. Most common, of course is a combinedthermal and mechanical internal boundary layer. Theconcept of a blending height has been introduced todenote the level, above which turbulent mixing haslevelled out horizontal gradients of climate variables.In the strictest sense, the study of microclimate is thenconfined to the airspace below the blending height.

The Surface Energy Balance

The way the surplus or deficit of radiative energy isredistributed or compensated for at the Earths surfaceis described through the energy balance equation,

which in a simplified form reads

NR HLvEG DS 1

Here, NR denotes the net radiation, Hand LvE are theturbulent fluxes of sensible and latent heat, respec-tively, and G is the ground heat flux. If the energybalance is set up for a volume instead of an infinites-imally thin surface such as the entire canopy layer, astorage term DSneeds to be included in eqn [1]. Thislatter term is the result of discernible divergence ofthe other energy fluxes involved and leads to warm-ing or cooling (drying or moistening) of the volume

considered.

The turbulent fluxes of sensible and latent heat arisefrom the fact that the atmospheric flow in the vicinityof the surface is generally turbulent. They constitutethe major mechanism of nonradiative vertical energytransport in the lowest portion of the atmosphere,

and are defined asH rcpw0Y

0

andLvE rLvw0

q0

,respectively, wherer is the air density, cp the specificheat at constant pressure,Y the potential temperature,Lv the latent heat of vaporization, and q the specifichumidity. The primed variables refer to the turbulentfluctuations with respect to a temporal average (rep-resented by an overbar).

The net radiation exhibits a strong seasonal anddiurnal cycle; typically, it assumes its maximumaround noon and is slightly negative during the night.In complex terrain, owing to orographic effects, theamplitude and phase of the net radiations daily cyclecan vary substantially over short distances, as can be

seen from Figure 3. In general, net radiation isdependent on the latitude (solar radiation), the surfaceproperties (reflectivity or albedo and emissivity), theatmospheric conditions (clouds) as well as the topo-graphy and surface cover of the surroundings (shadingeffects, longwave emission).

The partitioning of radiative energy surplus ordeficit into the other terms of eqn [1] again depends onthe properties of the underlying layer (vegetation, baresoil, rock, water, snow cover). In particular, theavailability of moisture at the surface or in the layerconsidered determines which of the latent heat flux orsensible turbulent heat transport is dominant. This is

most easily seen by considering the ratio between the

700

600

500

400

300

200

100

100

0

0 4 8 12 16 20 24

Time (LST)

NR(Wm

2)

Figure 3 Daily cycles of net radiation at four sites in an approximately northsouth-oriented alpine valley in southern Switzerland.

Shownare average data over 16 finedays with littleor no cloud in late summerfall1999.Opentriangles: vineyard site near thebottomof

the east-facing slope. Full dots: valley floor. Squares: meadow on the west-facing slope. Full triangles: alpine rubble site near the east-

facing ridge top.

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turbulent energy fluxes, the so-called Bowen ratio,expressed byb H=LvE.

Microclimate of a Vegetated Surface

Mean Vertical Structure

In the roughness sublayer, the state variables and thefluxes of energy, momentum, and mass all show apronouncedvertical variability. At a sufficient distancefrom the surface and under neutral conditions theprofile of mean wind speed can be shown to follow alogarithmic law:

uuz

un

1

k ln

zd

z0

2

whereun is the friction velocity as determined by the

Reynolds stress at the canopy top, k 0:4 the vonKarman constant,z0 the roughness length, and dthezero-plane displacement. A stability dependent mod-ification may be added to eqn [2] for nonneutralconditions. Both the zeroplane displacement and theroughness length are characteristic for the type andstatus of the underlying surface. They can be deter-mined from fitting observed wind speed profiles to eqn[2] or, alternatively, estimated from morphometricproperties of the surface. As a rule of thumb, d 0:7 hand z0 0:1 h can be used. Equation [2] is, strictlyspeaking, valid only in the inertial sublayer above theroughness sublayer. Within the roughness sublayer the

mean wind speed departs from the logarithmicbehavior, with a strong retardation at the mean heightof the roughness elements (Figure 4).

This retardation is due to the form and viscousdrags. Under certain simplifying assumptions themomentum balance equation for the roughness sub-

layer can be solved analytically to yield an exponentialwind profile:

uuz uuh exp a 1z

h

h i 3

wherea is a parameter dependent on the density andcharacter of the roughness elements and found to bein the range of 1 to 4 by fitting observed profiles toeqn [3].

Roughness sublayer profiles of the other climatevariables are also presented in Figure 4. The verticaldistribution of net radiation depends on the albedo ofthe canopy, as well as on the extinction of solarradiation (during daytime) and the absorption andemission of thermal radiation at different levels withinthe canopy. It is related to the distribution of biomassas given by the so-called leaf-area index. Owing to thegreater absorptance in the visible than in the near-

infrared range, there is a change in the spectralcomposition of the solar radiation flux from the topto the bottom of a stand.

The profiles of temperature and humidity andsimilarly CO2 (not shown) are complex. In foreststands they typically show an inflection point withinthe canopy. These profiles can be explained by thedivergence of the relevant energy and mass fluxes andby the distribution of the sources and sinks (includingthose at ground level), but do also reflect advectiveeffects. All profiles exhibit strong diurnal variations.

Turbulence statistics such as velocity variances orturbulent fluxes of sensible heat and momentum

exhibit a strong reduction from the canopy top downto the ground where very small or vanishing values areobserved. The most striking feature of canopy turbu-lence is probably the fact that it is governed by so-called coherent structures with spatial scales on theorder of the canopy height. As a consequence, the

2h

h

He

ight

0.0 0.5 1.0 1.5 2.0

u/uh

2 1 0.0015 0.0000 0.00150 1

h(K) qqh(kgkg1)

0 100 200 300 400

NR(Wm2)

Figure 4 Typical profiles of mean wind speedu, daytime potential temperature Y, specific humidityq, and net radiationNRwithin and

abovevegetated canopies.The lower portion of the specifichumidity profile is typical for moist surfacesonly. Forwind speedthe dashed line

corresponds to eqn [3] and the full line sketches a typical profile close to and within the roughness elements. Note that the profile for wind

speed qualitatively applies to both vegetated and urban surfaces.

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turbulent exchange of heat, moisture, or trace gaseswithin the canopy is often characterized by counter-gradient transport. This means that the so-calledK-theory, which is based on the assumption ofsmall dominating eddies and is a well-established

concept for the inertial sublayer, is not a usefuldescription of turbulent transport in the roughnesssublayer.

Owing to the rough nature of the surface, turbulentmixing is stronger in the upper part of the roughnesssublayer than in the inertial sublayer. As for the fluxesof sensible heat andspecific humidity (latent heat), thisincrease in most pronounced under near-neutral con-ditions. In this case the turbulent transport just abovethe canopy can be up to three times larger than in theinertial sublayer, whereas for momentum the enhance-ment is of the order of 10%.

Water and Carbon Budgets

There is an evidentfeedback between themicroclimateof a vegetatedsurface and vegetation itself. On the onehand, growth and composition of the canopy reflectthe boundary conditions set by the microclimate. Onthe other hand, vegetation exerts an active control onthe microclimatic conditions in its immediate sur-rounding, as seen when considering the water andcarbon budgets. Precipitation, interception, evapo-transpiration, and surface runoff, as well as infiltra-tion, drainage, and capillary rise, determine the waterstorage in the soil. Transpiration, however, is afunction of the water content and water potentialnot only in the soil but also in the roots, the stand andthe leaves, and depends therefore on the particularstate of the plants. Furthermore, photosynthesis andphotorespiration not only are regulated by physiolog-ical processes at the level of leaves and cells, but alsoare controlled by the overall radiation conditions andCO2concentration within the canopy.

In the case of agricultural surfaces, human activitiesmust be considered an important factor determiningthe microclimate of the stand. In fact, managementoptions are explicitly chosen to provide the bestgrowing and reproductive conditions for each partic-ular type of crop.

Microclimate of an Urban Surface

The microclimate of an urban surface differs from thatover an area covered by vegetation in many respects.The dominating roughness elements (buildings) arerigid and typically cover a considerably larger surfacefraction than is the case for, e.g., a forest. Consequent-ly, the retardation of the flowis even more pronouncedthan over a vegetated surface. Moreover, considerable

spatial variation in the thermal conditions can befound down to the level of the smallest entities, like thewalls of an urban street canyon. Further, the scale of anindividual roughness element is large enough for it tobe accessible for detailed experimental investigation.

The best-investigated entity of an urban surface istherefore the (generic) urban street canyon and itsimmediate surroundings.

The best-known characteristic of urban climate isthe so-called urban heat island. This refers to apronounced increase in near-surface temperature in acity as compared with its rural surroundings. Differ-ences are found to be strongest (several degreesCelsius) some hours after sunset and are attributedto, among others, the specific radiation conditionswithin block-like structures (street canyons), thedifferences in the turbulent diffusion of heat overurban and rural surfaces, respectively, and (at least in

winter) anthropogenic heat release.Effects of Building Geometry on Radiation,

Temperature, and Flow Characteristics

The daily cycle for the components of the radiationbalance differs for the various active surfaces within astreet canyon (sunlit and shadowed walls, streetcanyon floor). Long-wave incoming radiation is dom-inated by the emission from nearby walls, while short-wave (solar) incoming radiation is highly dependenton the orientation of the surface and may be severaltimes reflected on the various surfaces. Thereby, allradiation components can be shown to be related to

the so-called sky-view factor (i.e., the ratio of the skyseen at a particular point within the urban canopy tothat potentially available) and to the aspect ratio(width/height) of the canyon.

Within a street canyon the temperature may beslightly higher than aloft, owing to the trapping ofsolar radiation, but little spatial variation is observedexcept very close to the walls.

Profiles of mean wind speed over urban surfaces areusually described using the same concepts as forvegetated canopies (eqn [2]). As for the latter, the tworoughness parameters dand z0 are properties of theinertial sublayer and eqn [2] and, strictly speaking, are

valid only there.The transition from the inertial to the roughnesssublayer, and the actual profile within the roughnesssublayer, are strongly dependent on the orientation ofthe street canyon with respect to the ambient winddirection and its aspect ratio. For a narrow canyon theflow is often skimming, and wind speed almostvanishes within the canyon, leading to a stronggradient in wind speed around roof level. For widercanyons the exchange of momentum between thecanyon and the air aloft may be more efficient and the

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corresponding gradient in mean wind speed is some-what weaker. Generally, the magnitude of wind speeddoes not vary strongly with height within the lowerpart of the canyon, and wind speeds at the bottom ofthe canyon are found to be only a small fraction of

those at roof level (Figure 4).Coherent structures of the wind field (see above)are also found in urban environments. If the ambientflow is roughly perpendicular to the axis of a streetcanyon, a vortex can form inside the canyon(Figure 5), transporting ambient air downward alongthe downwind wall and canyon air upward along theupwind wall of the canyon. The strength and locationof this rotor is dependent on the canyon geometry andthe thermal conditions on the walls. For a streetwith significant traffic as a source of air pollutants,this flow pattern leads to distinctly different pollu-tant concentrations at the upwind and downwind side

of the street, respectively, if the street is alignedperpendicular to a dominant wind direction in theregion.

Air Pollutants and Turbulence Structure

Urban environments are characterized typically by ahigh density of sources for air pollutants near thesurface (traffic, domestic heating, and industry). Air

pollutant dispersion is the result of turbulent mixing,and this depends on the turbulence structure in thevicinity of the roughness elements. The height of theseelements is often less uniform than over vegetatedsurfaces, so that the average building height may notgenerally serve as a characteristic length scale. Rather,form drag is exerted on the flow up to the maximumheight of each roughness element. Consequently, thevertical turbulent transport of momentum (as the best-investigated measure for turbulent transport) is notapproximately constant (with height) as it is overideally homogeneous plant canopies or within theinertial sublayer. Rather, it assumes a maximum

magnitude at a level distinctly higher than the averagebuilding heightand decreases strongly as it approachesthe zeroplane displacement height.

Figure 5 Laser visualization of a vortex in a street canyon for perpendicularly approaching flow. (Reproduced with permission from

Kastner-Klein P (1999) Experimentelle Untersuchung der stromungsmechanischen Transportvorgange in Strassenschluchten. PhD

thesis, Institut fur Hydromechanik, University of Karlsruhe.)

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Below this level the momentum flux essentiallyvanishes, while velocityvariances (as a measure for theintensity of the turbulence field) have been reported tobe strongly dependent on the stability of the flow aloft.Owing to the relatively large dimensions of roughness

elements in a typical city, this layer of nonconstantstress can be several tens of meters deep (Figure 2).Model simulations suggest that an appropriate de-scription of the divergence of Reynolds stress in theurban roughness sublayer may have a large impact onthe dispersion of pollutants from low-level sourcesin cities.

Owing to its large roughness and the associatedproduction of mechanical turbulence, the stability ofan urban roughness sublayer is often found to benearly neutral during the night and only slightlyunstable during the day. Over an urban surface,conditions of extreme stability are therefore less likely

to be found than over nearby rural areas.

Nomenclature

cp specific heat at constant pressured zero plane displacementG ground heat fluxh average height of roughness elementsH turbulent (vertical) flux of sensible heatk von Karman constantLv latent heat of vaporizationLvE turbulent (vertical) flux of latent heatNR net radiation

q specific humidityuu mean wind speedun friction velocityw vertical component of the wind vectorz height coordinatezi boundary layer heightzn height of the roughness sublayerz0 roughness lengthb Bowen ratio

DS storage term in the energy balanceY potential temperaturer air density

See also

Boundary Layers: Overview. Downslope Winds.

LandAtmosphere Interactions:Overview. Mountain

Meteorology.Valley Winds.

Further Reading

Bailey WG, Oke TR and Rouse WR (1997) The SurfaceClimates of Canada. Montreal: McGillQueens Univer-sity Press.

Brutsaert WH (1982) Evaporation into the Atmosphere.

Dordrecht: Reidel.Campbell GS and Norman JM (1998) An Introduction to

Environmental Biophysics. New York: Springer.Cermak JE, Davenport AG, Plate EJ and Viegas DX (1995)

Wind Climates in Cities. NATO ASI Series E: AppliedSciences, vol. 27. Dordrecht: Kluwer.

Garratt JR (1992) The Atmospheric Boundary Layer.Cambridge: Cambridge University Press.

GeigerR, Aron R andTodhunter P (1996) TheClimate Nearthe Ground, 5th edn. Wiesbaden: Viehweg.

Hutchinson BA and Hicks BB (eds) (1985) The ForestAtmosphere Interaction. Dordrecht: Reidel.

Monteith JL (ed.) (1975)Vegetation and the Atmosphere,vol. 1, Principles. London: Academic Press.

Monteith JL (ed.) (1976)Vegetation and the Atmosphere,vol. 2, Case Studies. London: Academic Press.Monteith JL and Unsworth MH (1990) Principles of

Environmental Physics, 2nd edn. London: EdwardArnold.

OkeTR (1987)Boundary LayerClimates, 2ndedn. London:Methuen.

Rosenberg NJ, Blad BL and Verma SB (1983) Micro-climate: The Biological Environment, 2nd edn.New York: Wiley.

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